Heavy hydrogen in the stratosphere

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Heavy hydrogen in the stratosphere

T. Röckmann, T. S. Rhee, A. Engel

To cite this version:

T. Röckmann, T. S. Rhee, A. Engel. Heavy hydrogen in the stratosphere. Atmospheric Chemistry and Physics Discussions, European Geosciences Union, 2003, 3 (4), pp.3745-3768. �hal-00301153�

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3, 3745–3768, 2003 Heavy hydrogen in the stratosphere T. R ¨ockmann et al. Title Page Abstract Introduction Conclusions References Tables Figures J I J I Back Close

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Atmos. Chem. Phys. Discuss., 3, 3745–3768, 2003 www.atmos-chem-phys.org/acpd/3/3745/

Atmospheric Chemistry and Physics Discussions

Heavy hydrogen in the stratosphere

T. R ¨ockmann1, T. S. Rhee2, and A. Engel3

1

Max-Planck-Institut f ¨ur Kernphysik, Bereich Atmosph ¨arenphysik, Saupfercheckweg 1, 69117 Heidelberg, Germany

2

Max-Planck-Institut f ¨ur Chemie, Chemie der Atmosph ¨are, Becherweg 27, 55122 Mainz, Germany

3

Institut f ¨ur Meteorologie und Geophysik, Universit ¨at Frankfurt, Georg Voigt Str. 14, 60325 Frankfurt, Germany

Received: 25 June 2003 – Accepted: 15 July 2003 – Published: 25 July 2003 Correspondence to: T. R ¨ockmann (T.Roeckmann@mpi-hd.mpg.de)

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We report measurements of the deuterium content of molecular hydrogen (H₂) ob-tained from a suite of air samples that were collected during a stratospheric balloon flight between 12 and 33 km at 40◦N in October 2002. Strong deuterium enrichments of up to 400‰ versus Vienna Standard Mean Ocean Water (VSMOW) are observed, 5

while the H₂ mixing ratio remains virtually constant. Thus, as hydrogen is processed through the H₂reservoir, deuterium is accumulated in H₂. Using box model calculations we investigated the effects of H2sources and sinks on the stratospheric enrichments. Results show that considerable isotope enrichments in the production of H₂ from CH₄ must take place, i.e., deuterium is transferred preferentially to H₂ during the CH₄ oxi-10

dation sequence. This supports recent conclusions from tropospheric H₂isotope mea-surements which show that H₂produced photochemically from CH₄and non-methane hydrocarbons must be enriched in deuterium to balance the tropospheric hydrogen iso-tope budget. In the absence of further data on isoiso-tope fractionations in the individual reaction steps of the CH₄oxidation sequence, this effect cannot be investigated further 15

at present. Our measurements imply that molecular hydrogen has to be taken into account when the hydrogen isotope budget in the stratosphere is investigated.

1. Introduction

Molecular hydrogen (H₂), methane (CH₄) and water vapor (H₂O) are the three main hy-drogen reservoirs in the stratosphere. Once an air parcel has entered the stratosphere, 20

hydrogen can only be cycled between these species, since there are no net sources or sinks of hydrogen. Thus, the total hydrogen content χ ( ˆH₂)=2χ(CH4)+ χ(H2)+ χ(H2O), where χ denotes the mixing ratio, is generally constant in the stratosphere. Significant redistribution of total hydrogen can only occur during major dehydration events, which are very rare. During these events, ice crystals grow sufficiently large to fall to lower 25

altitudes, where they evaporate again.

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changes in its mixing ratio in the stratosphere up to 40 km (Ehhalt et al., 1977). It is not that H₂does not participate in the photochemical hydrogen cycling, but its production and loss rates are virtually identical. In the stratosphere, the most significant in situ source of H₂ is CH₄ oxidation (Fig. 1). Several reaction steps lead to the production of formaldehyde (HCHO), from which H₂ can be formed by photolysis. Figure 1 also 5

shows that at the most 2 of the 4 hydrogen atoms in a methane molecule can finally end up in H₂. One H atom is lost in the initial abstraction reaction, a second one in the reaction step CH₃O+ O2→ HCHO+ HO2. The fraction of the remaining two H atoms that form H₂through HCHO photolysis is dependent on the relative strengths of HCHO oxidation vs. photolysis and on the relative strength of the two photolysis channels. 10

The main stratospheric sinks of H₂are reaction with OH and O(1D) radicals. Reaction with Cl is a minor sink.

The end product of both the CH₄ and the H₂ oxidation chains is H₂O. This means that changes in atmospheric mixing ratios of both CH₄ and H₂have a potential impact on water vapor concentrations in the stratosphere. It is known that the tropospheric 15

increase in CH₄ mixing ratios, which is well documented (Blake and Rowland, 1988; Etheridge et al., 1992; Dlugokencky et al., 1998), has caused an increase in strato-spheric water levels (Oltmans and Hofmann, 1995; Engel et al., 1996). With H₂being projected as a major energy carrier in the future, emissions into the atmosphere during production, storage and transport of H₂are likely to increase (Tromp et al., 2003). This 20

could cause a substantial increase in stratospheric H₂O levels with severe implications for the energy balance of the earth (Forster and Shine, 2002), stratospheric tempera-tures (Forster and Shine, 2002), microphysical conditions in the stratosphere (Tromp et al., 2003) and stratospheric ozone levels (Evans et al., 1998).

Since the H₂ mixing ratio in the lower and middle stratosphere is nearly constant, 25

the net hydrogen cycling in the stratosphere can be regarded as a loss in methane and a production of water. Therefore, molecular hydrogen is not included in many studies that examine possible changes in the stratospheric hydrogen budget. How-ever, during the stratospheric processing of H₂ the isotopic composition may change

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although mixing ratios stay constant. This is particularly true for cycling between the three stratospheric hydrogen reservoirs, since their isotopic composition is very dis-tinct. Water vapor, on its way from the surface to the tropopause, loses heavy isotopes in condensation processes and becomes depleted in heavy isotopes. Several studies now indicate that water enters the stratosphere with an approximate isotopic compo-5

sition of δD ∼ −670‰ (Kuang et al., 2003). CH₄ and H₂ enter the stratosphere with their typical average tropospheric δ values of δD(CH₄) ∼ −86‰ (Quay et al., 1999) and δD(H₂) ∼ 130‰ (Gerst and Quay, 2000; Rahn et al., 2002b). Since both CH₄and H₂ are strongly enriched in D compared to H₂O, it is expected that H₂O formed via oxidation of these two gases will be enriched relative to the water that enters from the 10

troposphere. Therefore the deuterium content of stratospheric H₂O should increase as its concentration increases. Similarly, one might intuitively expect that H₂formed from CH₄should be isotopically light, because the CH₄is depleted in D relative to H₂. In ad-dition, a kinetic isotope effect in the CH4sink further depletes the CH₄that is removed (see below). On the other hand, a similarly strong fractionation in the removal of H₂ 15

by OH (HH is removed preferentially) enriches the remaining fraction of H₂. Additional isotope effects are expected in the oxidation pathway (Gerst and Quay, 2001). The net effect on the deuterium content of H2in the stratosphere is hard to estimate, because of the large differences in δ values between the hydrogen reservoirs and the kinetic fractionations involved. A first attempt to constrain the isotopic composition of strato-20

spheric H₂ from combined spectroscopic deuterium measurements on stratospheric water and methane (Irion et al., 1996), did not yield detailed information on δD(H₂).

In this paper we show that the heavy isotope content of molecular hydrogen is in-creasing with altitude in the stratosphere. The enrichment is surprisingly high with values reaching up to 400‰, and it correlates linearly with decreasing mixing ratios of 25

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Mixing and isotope ratios of hydrogen (and a suite of other trace gases like CH₄ and N₂O) were determined from whole air samples collected cryogenically during a strato-spheric balloon flight. Details of the sampler are given in Schmidt et al. (1987). Al-though H₂ does not condense at liquid Neon temperature, at which samples are col-5

lected, it enters the sampler entrained with the whole air flow and then cannot escape against the inflowing air under high flow conditions.

The deuterium content of H₂ is measured by continuous-flow isotope ratio mass spectrometry (CF-IRMS) using a method that has been recently developed. Details of the method will be presented elsewhere (Rhee et al. manuscript in preparation), and 10

here we describe it in brief. An aliquot of an air sample is condensed onto the cold head (∼40 K) of a liquid Helium compressor. H₂ does not condense at that temper-ature (neither do He and Ne) and is subsequently flushed with a slow flow of Helium onto a cryogenic trap filled with molecular sieve. The temperature of the liquid nitrogen coolant is reduced by pumping on the head space. When the sample has been col-15

lected, it is transferred to a cryo-focus trap immersed in liquid nitrogen at the head of a molecular sieve capillary gas chromatography column. The hydrogen is then released onto the column and admitted to the mass spectrometer via an open split interface. The reproducibility of the isotope ratio measurement is presently about ±3‰, as deter-mined from multiple measurements of a laboratory reference gas. The accuracy was 20

checked with commercial isotope standards (IsoTop, Messer Griesheim) with nominal isotope values of −9.5‰ and+205‰ and one reference gas whose isotopic ratio was determined by conventional dual inlet IRMS.

In addition to the D/H isotope ratio, H₂ mixing ratios can be readily obtained from the combined peak areas of the two isotopologues. Results show good agreement 25

with measurements carried out with a mercury oxide H₂ detector (T. Wetter, personal communication).

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expressed in δ notation as the relative deviation of the D/H ratio in a sample (SA) from a standard (ST), δ = ((D/H)SA/(D/H)_ST−1)∗1000‰. The international standard for hydrogen isotopes is Vienna Standard Mean Ocean Water (VSMOW) with an absolute D/H ratio of (155.76±0.05)·10−6(Hagemann et al., 1970). Kinetic fractionation factors in chemical reactions are expressed as the reaction rate of the heavy isotopologue 5

relative to the light isotopologue, e.g., α_HD−sink= k(HD−removal)/k(HH−removal).

3. Results

The altitude profile of the mixing ratio of H₂and its deuterium content determined from 13 air samples collected between 12 and 33 km over Aire sur l’Ardour, southern France (43.7◦N, 0.3◦W) on 24 October 2002 is shown in Fig. 2. It is known from previous 10

studies (Ehhalt et al., 1977), that the H₂ mixing ratio does not exhibit large variations throughout this altitude range. Its δD value, however, shows a pronounced increase from typical tropospheric values of about 130‰ at 12 km to nearly 400‰ at 32.4 km. Despite the fact that H₂ is produced in the stratosphere from isotopically much more depleted CH₄, it becomes actually strongly enriched. High stratospheric δD(H₂) values 15

have also been recently found by Rahn et al. (2002a).

Figure 3 shows δD(H₂) plotted versus the CH₄ mixing ratio, which is a proxy for the degree of photochemical processing in the stratosphere. It is evident that hydrogen gets progressively enriched in deuterium as CH₄ is destroyed further. The two pa-rameters show a very compact linear correlation (R2=0.998). We note, however, that 20

this does not necessarily imply a chemical connection. When we compare any two stratospheric species with local life times that are substantially longer than the trans-port times, their distribution is dominated by transtrans-port processes rather than chemistry, which does result in compact correlations (Plumb and Ko, 1992). We note that since this is true for both H₂ and CH₄ (Z ¨oger et al., 1999), the relation presented in Fig. 3 25

is not only characteristic for this single time and location, but is expected to hold (with possible small variations) throughout large regions of the stratosphere.

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The strong isotope enrichments convincingly illustrate that H₂ is not a mere spectator of stratospheric hydrogen cycling, but it plays an active role. Although the H₂ mixing ratio is constant, there must be a continuous production and destruction of H₂to cause the observed deuterium enrichment. If we assume that the general understanding of 5

hydrogen sources and sinks in the stratosphere is correct, this can only be due to either a faster production of HD compared to HH from CH₄oxidation, or a preferential destruction of HH, or both.

In their investigation of the tropospheric hydrogen budget, Gerst and Quay (2001) have investigated this issue in detail in an endeavor to explain the high δD value of H₂ 10

in the troposphere. We will discuss our stratospheric data along the same lines, but for stratospheric conditions. The fractionation in the stratospheric H₂sinks can be quanti-fied, since fractionation constants have been determined experimentally. The situation is less favorable for the stratospheric H₂source, i.e., production of H₂from CH₄. Frac-tionations of large magnitude are expected to occur in several reaction steps along the 15

reaction sequence (Fig. 1) (Gerst and Quay, 2001). Unfortunately, quantitative informa-tion is lacking for most of them, and it is not yet possible to model the transfer of deu-terium through the CH₄ oxidation chain. Therefore, at this stage we do not investigate the individual reaction steps and only attempt to answer the question: What isotopic composition is required for H₂ produced by CH₄oxidation to explain the stratospheric 20

observations? In the following, we name this quantity δD(H₂)_source since photochemi-cally produced H₂is the only molecular hydrogen source in the stratosphere. The aim is to determine a value for δD(H₂)_source which leads to a δD(H₂)−CH₄ correlation as shown in Fig. 3. A similar approach was also adopted from Gerst and Quay (2001). In their study, however, the situation was different due to the unknown relative strengths 25

of the photochemical and the soil sinks in the troposphere. In the stratosphere, only photochemistry is important.

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our simplified system. At the outset, we want to mention the limitations of using a box model for this purpose. They arise primarily due to the fact that a box model does not include effects of diffusion, transport and mixing. It has been shown, however, that these dynamical processes affect the isotopic composition of long-lived stratospheric trace gases, which are removed in the stratosphere (R ¨ockmann et al., 2001; Kaiser 5

et al., 2002). Generally, the ”apparent” fractionation constants which can be derived from stratospheric observations are significantly lower than the kinetic fractionation constants determined in the chemical removal reactions in the laboratory (Appendix A). This is taken into account by using the apparent fractionation constants rather than the actual kinetic fractionation constants in the box model. Whereas this approach leads 10

to realistic magnitudes of the isotope fractionation in box models, mixing is of course not treated realistically this way, because in reality mixing affects the isotope ratios by smoothing out gradients. This means that a modeled vertical profile may still include additional structure, e.g. a curvature, which is not sufficiently ”smeared out”, due to the inadequate way of including dynamical effects (in particular for the case of H2 which 15

also has a stratospheric source). Thus, reliable detailed altitude profiles can only be expected from at least 1D modeling with a realistic parameterization of vertical mixing. Knowing about these limitations, we perform the following box model calculations to illustrate the individual fractionation mechanisms and to put some constraints on the isotopic composition of H₂produced from CH₄oxidation.

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In our model, the starting point is always an air parcel entering the stratosphere with 1750 nmol/mol CH₄, 500 nmol/mol H₂and δD(H₂)=130‰. The isotopic composition of CH₄is irrelevant since we are only interested in the final product, i.e., δD(H₂)_source. At each model step, a small fraction of CH₄is removed, and H₂is produced with a certain yield y and isotopic composition δD(H₂)_source. To keep the H₂ mixing ratio constant, 25

the same amount of H₂is then removed again with the relevant fractionation constant α_HD−sink. Figure 4a illustrates the approach for a simple example. In the first box model run (green lines), H₂is produced with δD(H₂)_source=−80‰, which is the isotopic composition of tropospheric CH₄, the initial source material. Fractionations in H₂sinks

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are not included (αHD−sink=1), thus δD(H2)_source does not change as methane mixing ratios decrease. Light hydrogen is thus transferred to the H₂ reservoir and causes a considerable depletion although the mixing ratio is constant. The extent of hydrogen transfer depends on the fraction of hydrogen that is processed via H₂, and we assume two cases for the H₂ yield from CH₄ oxidation, y=0.6 and y=1.0, i.e., for each CH4 5

molecule destroyed 0.6 and 1 H₂ molecules are formed, respectively. As expected, when lesshydrogen is processed via H₂(y=0.6), the transfer is smaller.

In the second step, fractionations in the H₂ sinks are included. αHD−sink=0.76 is an estimate for a globally weighted average of the fractionations in the two sink reactions with OH and O(1D) (Appendix A), and αHD−sink,app=0.854 is the apparent fractionation 10

factor to be expected in the stratosphere under the influence of diffusive mixing (Ap-pendix A). The substantial isotope fractionation in the removal of H₂ does lead to an appreciable enrichment in the remaining H₂ fraction. Nevertheless, Figure 4a shows that the effect of the sink alone does not lead to the observed enrichments. Calcula-tions are shown again for a H₂ yield of y= 0.6 and y=1.0. As shown in Appendix A, 15

α_app is considered more realistic for the stratosphere (close to diffusion limited case), and we use this value as well as y=1.0 in the following.

The final parameter used in the model is the change of δD(H₂)_source with altitude. Since δD(H₂)_source does originate from CH₄, changes in the δD(CH₄) value lead to changes in δD(H₂)_source. The change of δD(CH₄) with CH₄mixing ratio in the strato-20

sphere can be predicted from the relative 13C and D fractionation constants for the sink reactions, observations of δ13C(CH₄) changes as a function of CH₄ mixing ratio and results from 2D modeling that are in good agreement with the observations for δ13C(CH₄). As shown in Appendix A, α_CH

3D−sink,app = 0.865 should be a realistic

ap-parent fractionation constant taking into account diffusive mixing, with the limitations 25

as discussed above.

Assuming initially a constant relationship between the CH₄ source material and the H₂ product in the stratosphere, we parameterize the change of δD(H₂)_source with de-creasing methane mixing ratio by the parameter α_CH

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creases with decreasing CH₄mixing ratio (Fig. 4b), which leads to higher stratospheric δD(H₂) values at smaller CH₄mixing ratios. Parametrizing the change of δD(H₂)_source with altitude by α_CH

3D−sink,appalso means that the choice of δD(H2)sourceis only free at

the tropopause, and we will denote this value δD(H₂)_s0. This is the value that we try to constrain in the following from our observations.

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The model is thus characterized by three parameters, namely α_{HD−sink,app}, α_CH

3D−sink,app, and δD(H2)s0, but clearly the values chosen above cannot explain the

stratospheric observations. Gerst and Quay (2001) concluded from tropospheric H₂ isotope budget considerations that δD(H₂)_source should be (130±70)‰ in the tropo-sphere to explain the high tropospheric δD(H₂) value. If δD(H₂)s0=130‰ is used in 10

the model, the resulting enrichments are closer to, but still lower than the observa-tions (Fig. 4c). Figure 4d shows that keeping the other parameters constant, the high stratospheric enrichments can be modeled fairly well using δD(H₂)_s0≈ 190 ± 40‰.

In the following, the other parameters of the model are varied. In Fig. 4e, α_{HD−sink,app} is changed from the value adopted above (0.854) in three steps to 0.84, 0.8 and fi-15

nally to the value in the reaction limit (0.76). This is a huge change, and most likely unrealistic, as argued in Appendix A. Comparing Fig. 4e to d shows that lowering the value of the parameter α_{HD−sink,app} has a similar effect as increasing δD(H2)_s0. Thus, when α_{HD−sink,app} is decreased sufficiently the model results are in the range of the observations already for δD(H₂)s0=130‰.

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In Fig. 4f the change in δD(H₂)_sourcewith altitude, parameterized by the value α_CH

3D−sink,app, is varied. Note again that changes in αCH3D−sink,app do not imply a

change in the fractionation associated with CH₄ oxidation, but only characterize the product H₂. Fractionations in other individual reaction steps in the CH₄ oxidation se-quence can cause an altitude dependence of δD(H₂)_sourcethat varies from the δD(CH₄) 25

profile. Two different values for αCH

3D−sink,app are chosen, 0.78, the value under

reac-tion limited condireac-tions and an even lower value of 0.68. δD(H₂)_s0 is adjusted again to yield results close to the measurements. In these runs parameters are chosen such that δD(H₂)_sourcevaries much more strongly with altitude (Fig. 4f). Interestingly, the

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vature in the modeled stratospheric profile disappears, which indicates that it strongly depends on the isotopic composition of the H₂ source. In fact, a good agreement of the modeled δD(H₂) profile and the data is obtained if the profiles of δD(H₂)_sourceand δD(H₂) are rather similar. As explained above, precise agreement between the altitude profile of our box model runs and the observations is not expected, due to the fact that 5

stratospheric mixing is only taken into account by adjusting the fractionation constants. In reality we expect mixing processes to remove some of the curvature seen in the calculated correlations.

Regarding the value of δD(H₂)_s0the model calculations show a quite consistent pic-ture. Including all the sensitivity tests, which cover a large range, values of δD(H₂)_s0 10

between 130‰ and 230‰ are required to reproduce the stratospheric observations. If we assume that δD(H₂)_s0, the deuterium content of H₂ produced from CH₄ near the tropopause is similar to δD(H₂)_sourcefor the troposphere, then this range can also be adopted for the troposphere. Thus, the stratospheric data constrain the range of (130±70)‰ for photochemically produced H₂ predicted by Gerst and Quay (2001) 15

based on tropospheric H₂ budget calculations to the upper half. A reduction of the range of the deuterium content of photochemically produced H₂by a factor of two puts a major constraint on the global H₂budget (Gerst and Quay, 2001).

The range of 130 to 230‰ for δD(H₂)_s0as derived in the box model runs indicates that an overall isotope enrichment of roughly 240 to 350‰ occurs in the oxidation se-20

quence from CH₄with δD(CH₄) ∼ −86‰ to the final H₂product (note that δ values do not add linearly). This massive enrichment must originate from one or more individual reaction steps in the CH₄oxidation sequence (Fig. 1). Unfortunately, little quantitative information about the individual reaction steps is available, which prevents a detailed investigation at present. As discussed in Gerst and Quay, (2001), at least in the initial 25

hydrogen abstraction step of the CH₄oxidation sequence (Fig. 1), H is expected to be removed preferentially, which would cause a deuterium enrichment in the final reaction product H₂. For more details the reader is referred to Gerst and Quay (2001), where all available isotope information about the reaction sequence is provided.

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High precision δD measurements on stratospheric H₂reveal a pronounced deuterium enrichment that increases with degree of photochemical processing in the strato-sphere. An approximately linear relationship between δD(H₂) and CH₄concentration is found with δD values increasing from 130‰ near the tropopause up to 400‰ at CH₄ 5

mixing ratios of 900 nmol/mol. The deuterium enrichments demonstrate that strato-spheric molecular hydrogen plays an important role in the stratostrato-spheric deuterium bud-get and has to be included in global budbud-get calculations. Box model calculations show that to explain the enrichment, H₂ produced from CH₄ oxidation must be strongly en-riched vs. the CH₄ source material, in agreement with conclusions from tropospheric 10

measurements (Gerst and Quay, 2001), and the value near the tropopause can be tightly constrained to δD(H₂)s0=(180±50)‰.

Appendix A: Fractionation constants in the stratosphere

The two major sinks of H₂in the stratosphere are reaction with OH and O(1D), with only a minor contribution from Cl (LeTexier et al., 1988). The relevant isotope effects have 15

been determined and are listed in Table 1. Whereas there is no kinetic fractionation in the reaction O(1D)+ H2, the reaction OH+ H2 proceeds almost twice as fast as OH + HD at typical stratospheric temperatures of 230 K. The global average stratospheric removal rate of H₂ by OH is similar to that by O(1D) (LeTexier et al., 1988). Thus the globally averaged kinetic fractionation factor is the average of the two individual 20

fractionation factors, thus α_HD−sink ≈ 0.76.

In the stratosphere, CH₄is removed by the three radicals OH, O(1D) and Cl. The rel-evant reactions have been characterized isotopically (Saueressig et al., 1996; Saueres-sig et al., 2001). Table 2 lists the fractionation constants for stratospheric temperatures. All oxidants preferentially remove light CH₄, and the remaining methane gets enriched. 25

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calculations of stratospheric CH₄isotope ratios (Bergamaschi et al., 1996; Saueressig et al., 2001). Since there are no high precision measurements of δD in CH₄ available yet, constraints for models thus far come from δ13C(CH₄). The model used by Sauer-essig et al. (2001) found a good match with observations (Sugawara et al., 1997) for global average removal rates of 41% oxidation by OH, 31% oxidation by O(1D) and 5

28% oxidation by Cl (C. Br ¨uhl, pers. comm.). A large set of new measurements (M. Braß et al., manuscript in preparation) confirms these numbers. The weighted global average fractionation constant for H abstraction from CH₄is thus α_CH

3D−sink = 0.78.

However, due to the effects of diffusion and mixing, the apparent (i.e. observed) fractionation factors α_app in the stratosphere are significantly smaller than the ones of 10

the removal reactions themselves (R ¨ockmann et al., 2001). This has been supported theoretically by Kaiser et al. (2002), who showed that α_appranges from α in the reaction limited case to√α in the diffusion limited case. Also, mixing of air masses with different isotopic composition decreases the apparent fractionations.

A comparison of reaction rate constants calculated by the 2D model (C. Br ¨uhl, pers. 15

comm.) and vertical eddy diffusion coefficients (Froidevaux and Yung, 1982) indicates that the stratospheric situation is in between the pure diffusion limited and reaction limited cases. Thus, to estimate suitable apparent fractionation constants we use the comparison of available laboratory fractionation data and stratospheric measure-ments. Comparison of laboratory and stratospheric isotope measurements for N₂O 20

(R ¨ockmann et al., 2001) demonstrate that we are close to the diffusion limited case in the lower and middle stratosphere. Also δ13C(CH₄) measurements show that the apparent fractionation factor α_13CH

4−sink,app ≈ 0.985 (Sugawara et al., 1997) is

signifi-cantly larger than α_13CH

4−sink≈ 0.975, the removal rate weighted fractionation constant

for the three sinks, and close topα_13CH

4−sink = 0.987, the value under diffusion

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ited conditions. In Table 3, our best estimates for stratospheric apparent fractionation constants in the removal reactions of CH₄ and H₂are shown. They are closer to the diffusion limited value, based on the observations of δ13C in CH₄. These numbers are used as starting values in the box model calculations presented in this paper. We note

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that, in the absence of experimental data, at this point we do not include in the model potential variations of the apparent fractionation factors with altitude, which may arise due to altitudinal variations in the removal rates of H₂and CH₄by the individual radical reactions.

Acknowledgement. We thank John Mak for help with the development of the analytical system.

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J ¨urgen Kiko provided the Helium compressor used in the analytical system. CH₄ concentra-tion measurements were obtained by Marc Braß, together with δ13C(CH₄) data (manuscript in preparation). Ingeborg Levin kindly provided her independent CH₄ concentration data which were used to validate our own CH₄concentration measurements. We are indebted to Christel Facklam and Ingeborg Levin for isotopic calibration of our mass spectrometer working standard

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H₂. Christoph Br ¨uhl provided his model results on stratospheric removal rates of CH₄ by the three individual sinks. We thank Jan Kaiser for valuable discussion about fractionation con-stants in the stratosphere and Jens-Uwe Grooß for a helpful discussion about box modelling. Konrad Mauersberger, Carl A. M. Brenninkmeijer and Jan Kaiser provided comments on the original manuscript.

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Etheridge, D. M., Pearman, G. I., and Fraser, P. J.: Changes in tropospheric methane between 1841 and 1978 from high accumulation-rate Antarctic ice core, Tellus, 44B, 181–294, 1992. Evans, S. J., Toumi, R., Harries, J. E., Chipperfield, M. P., and Russell, J. M.: Trends in

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H atom yield in O(1D)+ HD reaction, Chem. Phys. Lett., 253, 177–183, 1996.

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Table 1. Rate coefficients (in cm3molecule−1s−1) for the reaction of H2 and HD with O( 1

D) and OH for stratospherically relevant temperatures from (Talukdar et al., 1996; Talukdar and Ravishankara, 1996) and the corresponding fractionation factors

k(230 K) α(230 K)

O(1D)+ H₂ 1.1×10−10

O(1D)+ HD 1.1×10−10 1.00 OH+ H₂ 9.2×10−16

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Table 2. Fractionation constant for the CH₄ removal reactions from (Saueressig et al., 2001), globally averaged removal strengths for the different radicals derived from model calculations (C. Br ¨uhl, pers. comm.), and the calculated globally averaged fractionation constant

reactant share of total removal α(230 K)

OH 41% 0.735

O(1D) 32% 0.943

Cl 28% 0.626

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Table 3. Global average fractionation constants for the removal of CH₄ and H₂ in reaction limited conditions (α), diffusion limited conditions (√α), and the best estimate for a realistic α_appin the stratosphere

α √α α_app

H₂sink 0.76 0.872 0.854 CH₄sink 0.778 0.882 0.865

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Print Version Interactive Discussion © EGU 2003

C

OH, O(

1

_{D), Cl}

O

C

₂

O

₂

NO

C

O

₂

C

NO

₂ 2

CO +

hν

O

OH

C

OH, O(

1

_{D), Cl}

O

C

₂

O

₂

NO

C

O

₂

C

NO

₂ 2

CO +

hν

O

OH

H

₄

H

₃

H

HO,

H

,

H

Cl

H

₃

O

H

₃

O

H

₂

O

H

O

H

₂

H

CO +

H H

CO +

H

₂

H

₄

H

₃

H

HO,

H

,

H

Cl

H

₃

O

H

₃

O

H

₂

O

H

O

H

₂

H

CO +

H H

CO +

H

₂

Figure 1: Schematic diagram of the CH4 oxidation pathway that leads to H2 production. Out

of the 4 hydrogen atoms initially present (marked in red), two end up in H2CO and are

available for transfer into H2. The other two are transferred to other reaction products and

eventually end up as water. In particular in these H abstraction steps large fractionations can occur if light hydrogen is preferentially removed.

20

Fig. 1. Schematic diagram of the CH₄ oxidation pathway that leads to H₂ production. Out of the 4 hydrogen atoms initially present (marked in red), two end up in H₂CO and are available for transfer into H₂. The other two are transferred to other reaction products and eventually end up as water. In particular in these hydrogen abstraction steps large fractionations can occur if H is preferentially removed.

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Print Version Interactive Discussion © EGU 2003 450 500 550 600 10 12 14 16 18 20 22 24 26 28 30 32 34 A lti tu d e ( k m ) H₂ (nmol/mol) 100 150 200 250 300 350 400 450 10 12 14 16 18 20 22 24 26 28 30 32 34 δD(‰ )

Figure 2: Hydrogen mixing ratio (a) and δD(H2) in the stratosphere as a function of altitude.

Although the mixing ratio stays virtually constant, δD(H2) increases almost 300‰.

Fig. 2. Hydrogen mixing ratio and δD(H₂) in the stratosphere as a function of altitude. Although the mixing ratio stays virtually constant, δD(H₂) increases almost 300‰.

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Print Version Interactive Discussion © EGU 2003 800 1000 1200 1400 1600 1800 100 150 200 250 300 350 400 450 δ D ( ‰ ) CH₄ (nmol/mol)

Figure 3: δD(H2) plotted versus CH4 mixing ratio, which was independently determined on

the samples by Marc Brass [manuscript in preparation] with errors of ~1%.

22

Fig. 3. δD(H₂) plotted versus CH₄ mixing ratio, which was independently determined on the samples by Marc Braß (manuscript in preparation) with errors of ∼1%.

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Print Version Interactive Discussion © EGU 2003 -100 0 100 200 300 400 400 600 800 1000 1200 1400 1600 1800 a CH 4 (nmol/mol) data δD(H2)source δD(H2) αHD-sink=1 δD(H2) αHD-sink=0.76 δD(H2) αHD-sink=0.854 δD( H2 ) (‰ ) 400 600 800 1000 1200 1400 1600 1800 -100 0 100 200 300 400 b data δD(H2), αCH3D-sink=0.865 δD(H2)source, αCH3D-sink=0.865 CH4 (nmol/mol) δ D( H 2) (‰ ) 100 200 300 400 500 c δ D( H2 ) (‰ ) data δD(H2)source; δD(H2)s0=130‰ δD(H2); δD(H2)s0=130‰ 100 200 300 400 500 d data black lines: δD(H2)source

red lines: δD(H2) dotted: δD(H2)s0=150‰ solid: δD(H2)s0=190‰ dashed: δD(H2)s0=230‰ δD( H 2) (‰ ) 400 600 800 1000 1200 1400 1600 1800 100 200 300 400 500 e data δD(H2)source αHD-sink=0.76; δD(H2)s0=130‰ αHD-sink=0.80; δD(H2)s0=130‰ αHD-sink=0.84; δD(H2)s0=130‰ CH 4 (nmol/mol) δ D( H2 ) (‰ ) 400 600 800 1000 1200 1400 1600 1800 100 200 300 400 500 f data δD(H2); αCH3D-sink=0.78; δD(H2)s0=170‰ δD(H2); αCH3D-sink=0.68; δD(H2)s0=140‰ δD(H2)source; αCH3D-sink=0.78; δD(H2)s0=170‰ δD(H2)source; αCH3D-sink=0.68; δD(H2)s0=140‰ δ D( H 2) (‰ ) CH4 (nmol/mol)

Figure 4: Box model calculation results illustrating the effects of fractionation in hydrogen sources and sinks (a)-(c) and the sensitivity to the individual parameters (d)-(f). δD is plotted as a function of CH4 mixing ratio. a) includes only the fractionation in the H2 removal

reactions for three different fractionation constants and for H2 yields from CH4 of y = 1.0

23 Fig. 4. Caption on next page. For detail use zoom tool.

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Fig. 4. Box model calculation results illustrating the effects of fractionation in hydrogen sources

and sinks(a)–(c) and the sensitivity to the individual parameters (d)–(f). δD is plotted as a

function of CH₄ mixing ratio. (a) includes only the fractionation in the H₂ removal reactions for three different fractionation constants α_HD−sink and for H₂ yields from CH₄of y=1.0 (molec H₂/molec CH₄removed), solid lines, and y=0.6 (molec H₂/molec CH₄removed), dotted lines. In (b) the change of δD(H₂)_source with CH₄ mixing ratio, parameterized by α_CH

3D−sink,app = 0.865,

is added. This leads to an enrichment of δD(H₂)_source with decreasing concentration. In (c)

δD(H₂)_s0=130‰ from Gerst and Quay is used, but even higher values of δD(H₂)_s0=150‰ to 230‰ (d) are necessary to bring the model results to the range of the observations. In (e) the fractionation factor for the H₂sink, α_{HD−sink,app}is varied, in (f) α_CH

3D−sink,app is varied. δD(H2)s0